, pilar llanes2,3*, eli silver2, gary hoffmann , steve ...ward/papers/ritter-deposits.pdf · 20...

Submarine landslide deposits as a key to understanding tsunami generation by volcano lateral 1

collapse: The historical collapse of Ritter Island, Papua New Guinea 2

3

Simon Day1, Pilar Llanes2,3*, Eli Silver2, Gary Hoffmann4, Steve Ward2, and Neal Driscoll5 4

5

1Dept. of Earth Sciences, University College London, London, UK 6

2Earth and Planetary Sciences Dept., University of California, Santa Cruz, CA 95064, United 7

States 8

3Dept. Geodinámica, Geológicas, Universidad Complutense de Madrid, Madrid 28040, Spain 9

4CFD Lab, Mech. Engineering, University of California, Santa Barbara, CA 93106, United States 10

5Scripps Institution of Oceanography, La Jolla, CA 92093, United States 11

*Corresponding author 12

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ABSTRACT 14

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The March 13th 1888 collapse of Ritter Island in Papua New Guinea is the largest known sector 16

collapse of an island volcano in historical times. One single event removed most of the island and 17

its western submarine flank, and produced a landslide deposit that extends at least 70 km from the 18

headwall of the collapse scar. We have mapped and described the deposits of the debris avalanche 19

left by the collapse using full-coverage bathymetry, side-scan sonar images, seismic-reflection 20

profiles, TowCam photographs of the sea bottom and scarce samples from a dredge. Applying 21

concepts originally developed on the 1980 Mount St. Helens collapse landslide deposits, we find 22

that the Ritter landslide deposits shows three distinct morphological facies: large block, matrix-23

rich and debris flow distal facies. Restoring the island’s land and submarine topography we 24

obtained a volume of 4.2 km3 for the initial collapse, about 75% of which is now forming the 25

large block facies at distances less than 12 km from the collapse scar. The matrix-rich facies 26

volume is unknown, but large scale erosion of the marine sediment substrate yielded a minimum 27

total volume of 6.4 km3 in the distal debris flow and/or turbidites deposits, highlighting the 28

efficiency of substrate erosion during the later history of the landslide movement. Our analysis 29

shows that well-exposed submarine landslide deposits can be interpreted in a similar way to 30

subaerial volcano collapse deposits, and that they can be useful to interpretate older, incompletely 31

exposed landslide deposits. Studying the deposits from a facies perspective provides the basis for 32

understanding the mechanisms involved in the movement and emplacement of deposits from a 33

collapse event, and for reconstruction of its kinematics, both of which are crucial for improving 34

tsunami modeling. 35

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1. Volcano lateral collapse and tsunamis: the study of submarine landslide deposits 37

38

Lateral collapses are a near-ubiquitous feature of volcanoes in most island volcanic arcs. 39

Numerous collapse generated landslide deposits with volumes of cubic kilometers to tens of cubic 40

kilometers occur both in volcanic island arcs such as the Lesser Antilles (Deplus et al., 2001; Le 41

Friant et al. 2003) and Kermadec arcs (Wright et al., 2006), and in back arc settings such as the 42

Sea of Japan (Satake & Kato 2001). Island or coastal volcano lateral collapses with volumes of a 43

few cubic kilometers and maximum thicknesses of 500 meters or more occur at a global 44

frequency of the order of ~ 1 per 100 years (REFERENCES?). Smaller landslides removing 45

summit dome complexes or thin layers from the volcano flanks occur at intervals of years to 46

decades (e.g., 2002 Stromboli; Bonaccorso et al., 2003). 47

Ritter Island 1888 is the largest ocean-entering volcano lateral collapse, in volcanic arc 48

settings, for which we have written accounts of both the collapses and the resulting tsunamis 49

(Table 1). It produced a large, regionally destructive tsunami that was witnessed by literate 50

observers who timed the waves with watches and provided the most detailed accounts of any 51

tsunami produced by a km-volume lateral collapse into the ocean (summarized in Cooke, 1981). 52

That data has been key to modeling the generation of the tsunami by the landslide and its later 53

propagation (Ward & Day 2003). Understanding how submarine landslides from volcano 54

collapses generate these tsunamis is a key motivator for studying their kinematics and ultimately 55

their mechanics. There is a need for detailed studies of young, well exposed landslide deposits on 56

the ocean floor that are associated with known tsunamis in order to provide a basis for the 57

interpretation and modeling of older, incompletely exposed landslide deposits from other island 58

arc volcanoes. 59

In this paper we describe the submarine landslide deposit from the historical collapse of 60

Ritter Island in Papua New Guinea (Figure 1, Figure 2). We show that whilst the collapse was a 61

single event, the deposit is subdivided into distinct facies as a result of rheological 62

transformations during movement; and that the geometry of the deposit and associated erosional 63

features can be used to reconstruct the kinematics of landslide movement. We need detailed 64

studies of young, well exposed volcano collapse landslide deposits on the ocean floor in order to 65

provide a basis for the interpretation of older, incompletely exposed landslide deposits. Historical 66

descriptions, sketches and measurements of pre-collapse Ritter Island define its geometry above 67

sea level in more detail than for any other volcano lateral collapse, and these descriptions also 68

constrain the sequence of events (Johnson et al.,1972). These data have provided us with a unique 69

opportunity to analyze the landslide facies of a submarine volcano collapse event. 70

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2. Background 72

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2.1 Field observations from Ritter Island 74

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Descriptions by Johnson et al. (1972) and our onshore field observations in 2006 reveals that the 76

subaerial exposed part of Ritter is a volcanic sequence dominated by rubbly or spatter-fed lavas 77

with some massive flow cores, and intercalations of finer-grained scoria. Steeply east dipping 78

sequences form the whole of the volcanic stratigraphy, from the rim of the collapse scar to the 79

oldest subaerial lavas, exposed adjacent to hyaloclastite tuffs with thin scoria beds that are the 80

structurally lowest and oldest unit of the remaining subaerial part of the edifice. These rocks are 81

cut by numerous irregular, broadly N-S striking dikes and outward dipping sheet intrusions up to 82

several meters thick, and some NW striking dikes. Fumarolic alteration is present in the lower 83

part of the exposed sequence. The island is incised by channels or gullies up to several tens of 84

meters wide, with steep to vertical walls up to a few tens of meters high. Similar gullies filled by 85

lava flows and discrete boulder brecciate units form minor unconformities within the pre-collapse 86

volcanic sequence. However, there are no marked unconformities extending along the face of the 87

collapse scar that would record past collapse events, nor are there flat lying, ponded lava flows 88

that would indicate a collapse scar buried within the intact remnant of the island, as seen 89

elsewhere (Day et al. 1999; Carracedo et al., 1999; Carracedo et al., 2002). There is, therefore, no 90

evidence from the island for recent pre-1888 lateral collapses of Ritter, comparable to that which 91

occurred in 1888. On the other hand, the pre-1888 sketches of the island (Figure 2) show evidence 92

for smaller landslide scars that would have been produced in events that removed rock volumes of 93

less than 0.1 km3. Comparison with data for similar events such as the 2002 Stromboli landslide 94

(Table 1B) suggests that deposits from such landslides would not have gone more than a few 95

kilometers from Ritter and would therefore be either buried under the 1888 deposits or would 96

have been eroded and entrained into the 1888 debris avalanche. 97

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2.2. Historical observations from Ritter Island 99

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Cooke (1981) summarized the observations of activity at Ritter Island from William 101

Dampier’s sighting of an eruption in 1700 through to the German colonization of northern New 102

Guinea that started in the 1880s. These accounts indicate that Ritter experienced intermittent 103

eruptive activity in the years just before the collapse, but numerous accounts of fumaroles indicate 104

hydrothermal activity. Later collection of reports from communities on Umboi Island indicated 105

that the collapse of the volcano around 6 am on March 13th 1888 was preceded by explosions 106

accompanied by earthquakes, but there was no mention of eruptions of magma (Fisher, 1939; 107

1957). There is no evidence for a large magmatic eruption associated with the collapse, although 108

it seems likely that at least hydrothermal or phreatic explosions took place. Thus, the collapse is 109

most likely to have been associated with phreatic but not magmatic explosive activity. This 110

activity is of the Bandai type, following the classification of Siebert et al (1987). 111

The collapse produced tsunami waves that were recorded at observation points between 112

100 and 600 km from the volcano between about 6:30 am and 8:30 am local time on March 13th, 113

and continued for 40-150 minutes with dominant wave periods around 3 minutes. No observer 114

described more than one series of tsunami waves. Ward & Day (2003) showed that the times of 115

arrival and durations of the tsunami at these different sites are consistent with the propagation and 116

progressive frequency dispersion of ~3 minute period tsunami waves produced by a single 117

landslide from Ritter Island. No further tsunamis were recorded at the various observation points 118

in the days, weeks and months immediately after the event. The historic evidence is consistent 119

with only one major collapse or landslide event occurring in 1888. 120

Earthquakes and submarine explosive activity immediately west of Ritter Island in 1972 121

and 1974 marked the growth of a new submarine volcanic vent within the collapse scar (Cooke et 122

al. 1976; Cooke, 1981). Two small tsunamis occurred in these years, the only ones to be recorded 123

from Ritter between 1888 and the time of our survey in 2004. Any young, large volume landslide 124

deposits in the area north west of Ritter Island are most likely associated with the single 1888 125

collapse. 126

Despite the lack of large volume tephra falls associated with the events of 1888, it was 127

supposed for many years that an explosive eruption and caldera collapse had occurred at Ritter 128

Island in 1888 (Fisher, 1957; Cooke, 1981). Following the 1980 May 18th Mount St. Helens 129

lateral collapse, the Ritter event was investigated by Johnson (1987) with a single-beam 130

echosounder survey of the immediate area of the island that revealed a broadly west facing open 131

horseshoe shaped collapse scar and large bathymetric highs interpreted by Johnson as landslide 132

blocks, allowing him to estimate a volume for the collapse scar of 4- 5 km3. The later survey did 133

not extend far beyond the collapse scar and so the full run-out distance, extent and structure of the 134

deposits associated with the collapse has been an important gap in knowledge of the 1888 Ritter 135

Island event. 136

137

2.3. Restoration of pre-collapse island topography 138

139

Observers of Ritter in 1700 described the island as a steep, gullied cone with a height 140

about half of the adjacent extinct volcano of Sakar Island rising at 998 m (Dampier, 1729). 141

Accounts in the 19th Century emphasized the extremely steep-sided nature of the subaerial part of 142

Ritter volcano with measurements of about 780 m height and 1.83 km E-W width (Dumont 143

d’Urville, 1832), implying slopes significantly in excess of 45°. To estimate the island’s 144

dimensions we compared the sketch and description of the island, made in early 1835 by Jacobs 145

(1844), with the modern island. Matching the present island’s features to gullies in Jacobs’ sketch 146

(Figure 2). Scaling the sketch to the modern island with a maximum height 140 m yields a pre-147

collapse height of around 750 m and an E-W width of around 1.5 km. These pre-collapse 148

dimensions are consistent with the ~45° slope of the surviving eastern flank of the island. Because 149

the island is still about 1.9 km long it must have been elongated N-S. Another sketch of the island 150

made on the same voyage in 1835 by Woodworth (Figure 2) confirms the overall steep slopes of 151

the island. Both sketches suggest that the western side of the island (to the left in both cases) was 152

steeper and perhaps had already experiencing landslides. Our bathymetric data, discussed below, 153

indicates that the submarine slopes of the volcano down to 700 m below sea level are mainly 154

symmetrical and an upward extrapolation of these slopes to sea level implies that the pre-collapse 155

island had a N-S length of a little over 2 km. We use this restoration of the pre-landslide 156

topography together with an extrapolation of the submarine collapse scar in a later section to 157

approximate the volume of the sector of the island collapsed during the 1888 event. 158

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3. Methods 160

161

During November-December of 2004 we conducted a geophysical research cruise along 162

the Bismarck volcanic arc aboard the R/V Kilo Moana. We acquired swath bathymetry data with 163

a Simrad EM-120 hull-mounted echo sounder, operating at 12 kHz (±0.75 kHz). The beams are 164

transmitted across a total swath opening of up to 150°, depending on depth. Acoustic backscatter 165

intensity images were obtained with the HAWAII MR1 Sonar, a portable shallow-towed system 166

with swath width up to 7.5 times the water depth. The MR1 sonar transducers are housed in a 4.8 167

m long, 1600 kg vehicle that is towed beneath the surface mixed layer (80 to 120 m) and requires 168

a minimum water depth of 500 m, limiting our work at shallower depths. The Hawaiian Mapping 169

Research Group processed both the bathymetry data and the acoustic images on board using 170

Simrad’s software. We used a median filter and gridded the bathymetry and the acoustic data to 171

24 and 16 m, respectively. High-resolution Compressed High Intensity “Radar” Pulse (CHIRP) 172

sub-bottom sonar data were collected with an X-Star towed 1-5.5 kHz instrument (Scripps 173

Institution of Oceanography). CHIRP data were acquired by transmitting approximately every 5-174

10 m along selected transects with penetration up to several tens of meters. CHIRP data are 175

unmigrated. 176

Digital photographs of the seafloor were taken using the digital “TowCam” system 177

(Woods Hole Oceanographic Institution) along four transects, over the Ritter deposits, each of the 178

tows lasting approximately 4 h from deployment to recovery (see Figure 3 for locations). The 179

camera is towed 100-300 meters behind the ship at speeds of ¼ to ½ knot, at an elevation of 5 180

meters above the seafloor. The TowCam records depth, altitude, water temperature, water clarity 181

or turbidity and it was set to record overlapping seabed photos every 10 s. It allowed us to 182

document the location and distribution of a variety of different lithologies and structures exposed 183

on the seabed, and distribution of megafauna (animals visible in camera were typically >1 cm). 184

Correlation of the camera images with the bathymetric, acoustic backscatter and CHIRP data (all 185

of which were used to plan the TowCam lines), has provided ground truth and therefore allowed 186

better interpretation of the various types of acoustic data. 187

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4. The submarine collapse scar and the post-collapse volcanic edifice 189

190

The collapse scar has a 2 km long headwall in a N-S direction and extends to the western 191

foot of the volcano at 900 m depth, where it is up to 4 km wide with the sidewalls defining a 192

sector arc of about 60 degrees (Figure 4). The north and south walls of the collapse scar are steep, 193

rising 100 to 200 m above the level of the post-collapse fill, and with very high backscatter 194

intensity suggesting extensive exposure of bedrock on the cliff faces (Figure 5). The spoon shaped 195

geometry of this surface may suggest that the landslide began as a single rotational block failure. 196

The single continuous tsunami reported by the eyewitnesses indicates that the largest part of the 197

collapse scar was generated in a single event; but the failure may have consisted of a number of 198

blocks moving in rapid succession upon a single failure surface, as seen directly in the 1980 199

lateral collapse of Mount St. Helens (Voight, 1981; Voight et al., 1983), and thereafter moved as a 200

single unit. No lateral boundaries between separate lobes of similar facies deposits are seen in the 201

Ritter Island deposits. If multiple collapse events had occurred, deposits from the earlier collapse 202

events would be expected to show a partial cover from deposits associated with the later events 203

(Yokose & Lipman, 2004). We see no evidence to indicate that the collapse scar was produced by 204

multiple failure events. 205

The collapse scar has been partly filled by a post collapse volcano with a summit crater a 206

few hundred meters across and a rim at about 300 m depth. This and other volcanic vents on the 207

surface of the landslide deposit are distinguishable from collapse blocks because they have 208

conical shapes and low backscatter intensity (V in Figure 5). The summit vent of the large edifice 209

within the collapse scar is located at the sites of eruption plumes that broke the surface of the sea 210

during the 1972 and 1974 explosive volcanic eruptions (Cooke et al., 1976) dredged by the later 211

authors and found to be XXXX. Eruptive activity inside the scar left by a lateral collapse is 212

common because ascending magmas take advantage of the weakened areas created by collapse 213

structures, erupting within it (e.g. the Lesser Antilles, Boudon et al., 2007). The flanks of the main 214

post-collapse volcano are smooth, lacking discrete flank vents, and have uniform moderate 215

backscatter intensity. The smooth lobes of low backscatter intensity extending from the base of 216

the post collapse edifice to the NW and SW may also be features of post collapse eruptions, such 217

as deposits from syneruptive volcaniclastic density currents. 218

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5. Submarine landslide facies 220

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Glicken (1991; 1996) subdivided the morphological and sedimentological characteristics 222

of the deposits from a volcano collapse into distinct facies. Landslide facies for the Mount St. 223

Helens deposit were described in great detail, including block size distribution, abundances of 224

discrete blocks in contrast with matrix material, deposit thickness variations and depositional 225

slopes. To gain a more complete understanding of emplacement, we compared Glicken’s 226

landslide facies concepts to our submarine observations. The use of full-coverage bathymetry, 227

backscatter images, and TowCam photographs of the sea bottom, samples from cores and dredges 228

and high-resolution seismic data, has allowed us to define and map the submarine landslide facies 229

of the 1888 Ritter Island collapse. We have divided the deposit into three facies: large block 230

avalanche facies, matrix-rich debris avalanche facies, and distal debris flow facies. Such a 231

approach will yield new insights into how submarine avalanche deposits evolve downslope away 232

from their source. 233

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5.1. Large block debris avalanche facies 235

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5.1.1. Depositional features 237

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The large block facies extends 12 km from the headwall of the collapse scar, filling the 239

full width of the valley between Sakar and Umboi islands as far as a constriction in the valley 240

formed by two overlapping ridges trending NE-SW across the valley sides (R1 and R2, Figure 4). 241

The margins and top surface of the large block facies deposit are at a depth of about 900 m on its 242

eastern and northern sides, but descend towards the gap between the ridges, where the top of the 243

deposit is at 1240 m depth at its furthest downslope limit. The average slope along the length of 244

this facies is about 2.1°. This area has large discrete irregular blocks or hummocks up to more 245

than 1 km across and over 100 m high, with generally high backscatter intensity (B in Figure 5). 246

The distribution was determined by contrasting acoustic responses of the lava blocks and the fine 247

products surrounding them, as well as the angular shapes of the blocks, which provide surfaces 248

facing the beam geometry. The largest of the blocks (B in Figure 6) is located in the mouth of the 249

collapse scar and was first identified by Johnson (1987). Its asymmetry is consistent with it being 250

a displaced and back-rotated segment of the flank of the volcano. The majority of the other large 251

blocks are concentrated toward the southern and western margins of the deposit. The high slopes 252

and irregular surfaces of the blocks produced strong seabed diffractions and the characteristic lack 253

of seismic penetration over blocks deposits. The absence of a sedimentary cover over these large 254

blocks in the seismic profiles (Figure 7) is consistent with the recent age of the landslide deposit. 255

The overall geometry of the large block facies deposit suggests that it was ponded in the 256

valley to the east of the constriction. This ponding is similar to that seen in the large block facies 257

of the Mount St. Helens landslide deposit (Glicken 1991, 1996) and others like Mount Shasta 258

(Crandall et al. 1984, Crandall 1988). 259

260

5.1.2. Erosional features 261

262

On the lower slope of Umboi Island there is a laterally extensive region of uniform high 263

backscatter intensity with a sharp upslope contact (“Scouring” in Figure 5) sub-parallel to the 900 264

and 800 bathymetric contours. This contact undulates, curving upslope and over a small ridge (R3 265

in Figure 4) and reaching its maximum elevation at a depth of 700 m, which is 200 m above the 266

level of the large block facies deposit (R1 in Figure 4). The contact is continuous across several 267

fans of high backscatter material trending down the slope from Umboi Island, but it is 268

discontinuous where larger fans are present. The high backscatter zone also extends over the crest 269

of the larger ridge (R1 in Figure 4) that bounds the large block facies deposit at the constriction 270

and down slope into the matrix rich facies deposit area. 271

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5.2. Matrix rich debris avalanche facies 273

274

The transition from block-rich to matrix-rich facies occurs where the gap between Umboi 275

and Sakar is constricted between two ridges (R1 and R2 on Figure 4). Similarly, the transition 276

from large block facies to matrix-rich facies in the 1980 Mount St. Helens debris avalanche 277

deposit occurs at a constriction in the North Fork Toutle River (Glicken 1991, 1996). In this 278

section we describe the nature of depositional and erosional features, the latter consisting of steep-279

sided mesas and grooved terrain. 280

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5.2.1. Depositional features 282

283

The matrix rich facies is characterized by a distinctive speckled pattern in the backscatter 284

intensity map (Figure 8), and as a coarse blocky and non-stratified mass in high-resolution 285

seismic data (Shots 2800-3500 in Line 2/7, shots 700-1300 in Line 3/1 Figure 9) that corresponds 286

to irregular hummocky terrains, which is typical of submarine debris avalanche deposits (Moore 287

et al., 1989, 1994; Watts and Masson, 1995). We interpret high reflectivity patches to generally 288

indicate the presence of blocks. High reflectivity results from a high contrast in the sound velocity 289

times the density of the materials (impedance contrast). Volcanic blocks have high impedance 290

contrast relative to the sediment deposits surrounding them. The patches of high backscatter 291

intensity in the upper area of the rich matrix facies are 100-300 m across. 292

The acoustic character, a lack of CHIRP penetration, and discrete areas of diffractions 293

allowed us to map the lateral extent of the matrix rich facies, differentiating it from the sediments 294

of the lower slopes of Umboi and Sakar islands. Along the lower slope of the islands, well-295

layered reflectors are observed. Diffractions are present in areas of high slope. The thickness of 296

acoustically laminated sediments varies depending on slope and location relative to the islands. 297

Line 3/1 in Figure 9 shows as much as 0.016 two way time seconds of sediments along the slope 298

of Umboi and 0.008 two way time seconds of sediments along the slope of Sakar. This acoustic 299

penetration is equivalent to 12 and 6 m using a nominal water velocity of 1500 m/s to convert for 300

thickness. 301

A series of lobes occurs at the downslope end of the matrix rich facies deposit. These 302

lobes are located in flat-floored channels between a series of steep-sided, flat-topped features 303

referred to here as mesas. TowCam images of these distal lobes with their characteristic mottled 304

backscatter intensity show that the lobes are partly covered in fine-grained, bioturbated (hence 305

soft and unconsolidated) marine sediment. Hummocks encountered along the TowCam lines 5 306

and 6 (Figure 3) through the lobes are a few tens of meters to several tens of meters across, and 307

several meters high, composed of poorly sorted, angular blocks mainly less than a meter across, 308

but ranging in size from centimeters to over five meters long (Figure 10a, b and c). No banding or 309

layering is observed in the blocks, nor are there any clear variations in block characteristics at any 310

one hummock or mound, suggesting that these are monolithologic, whereas different mounds 311

show different rock types. Polygonal-jointed columnar blocks from a thick jointed substrate are 312

exposed in one of the mounds of TowCam 5 (Figure 10b), whereas other mounds in the same 313

TowCam line showed pitted clasts that appear to be coarsely vesicular (Figure 10c). The clasts 314

form a habitat for abundant sessile filter-feeding organisms like white sponges, soft corals, 315

anemones and crinoids that require hard substrates for their attachments and are not observed in 316

sediment-draped areas, except for rare examples of anemones attached to sunken tree branches. 317

We interpret the mounds as debris avalanche hummocks or fragmented blocks composed of 318

volcanic rocks from the Ritter Island edifice. 319

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5.2.2.1. Submarine canyon 323

324

Along most of the length of the matrix rich facies deposits there is a 1 km wide and 20-40 325

m-deep submarine canyon. It starts at the valley constriction that limits the block and matrix 326

facies deposits and it is easily recognized out to the 1700 m isobath. The five CHIRP profiles 327

crossing the channel image a strong reflection with no penetration, indicating the lack of a 328

sedimentary sequence in the canyon (e.g., line 3/1 in Figure 9) and demonstrating that this canyon 329

is erosional. 330

331

5.2.2.2. Black Streak 332

333

Between the distal lobes of the debris avalanche deposit there is an elongated feature with 334

very high backscatter intensity that we named the “Black Streak” (BS in Figure 8). It is 1 km wide 335

and rises 20 m above the surrounding terrain. Images over its crest show that it is composed of 336

angular clasts that are dark and pitted, appearing vesicular or scoriaceous (Figure 10d). An 337

abundant and diverse fauna of sponges, soft corals and crinoids is present, similar to that seen on 338

the debris avalanche blocks. Unlike the poorly sorted clasts in the debris avalanche deposit, the 339

clasts in the “black streak” are uniform in size, 10 to 20 cm across. The CHIRP profile crossing it 340

shows its relief with strong echoes underlying transparent facies and overlying a debris avalanche 341

deposit that covers its eastern side (Line 24, Figure 11). 342

343

5.2.2.3. Mesas 344

345

To either side of the lobes of the matrix-rich debris avalanche deposit there are several 346

flat-topped surface features rising above the surrounding seabed and with a mesa-like appearance 347

(M in Figure 8). These mesas are unique to this area, and similar features were not found 348

anywhere else in the length of the Bismarck arc covered by the research cruise. The best-defined 349

example (M* in Figure 8) has a near vertical scarp, 50 m-high, with a steep slope covered by 350

cohesive sediment blocks. Along its margins, except the NE one, there are marked troughs, 351

several m-deep and almost 100 m-wide, where exposed lithic clasts similar to the clasts seen in 352

the matrix-rich debris avalanche deposit were seen across it in TowCam 5. Two of those mesas 353

were also crossed by CHIRP profile 24 (Figure 11), revealing the presence of sedimentary 354

layering and differentiating them from the nearby avalanche deposits. 355

Along the SW margin of the matrix rich facies deposit, there are several mesas developed 356

on a marked slope above the debris avalanche deposit and without troughs developed around 357

them (M’ in Figure 8; Line 3/1 in Figure 9). 358

359

5.2.2.4. Grooved terrain 360

361

To the N and E of the matrix rich facies deposits and on the tops of the sediment mesas, 362

we observed a lineated or grooved terrain in the backscatter and TowCam images (mapped with 363

dotted areas in Figure 12). A correlation of sets of TowCam images and the camera altimeter 364

allowed us to estimate that the largest grooves imaged were several tens of meters across and 365

several meters deep and seem to be small-scale counterparts of larger linear structures (several 366

thousand meters long) that produce the fabric seen in the backscatter data. The walls of the 367

grooves are steep, occasionally overhanging, and characteristically scalloped suggesting that the 368

grooves are cut into cohesive sediment with significant strength (Figure 10e, f). In rare cases, 369

planar sedimentary bedding is visible in groove walls. A sessile fauna dominated by near-370

transparent barrel-shaped sponges occurs on some groove walls but the diverse fauna of sponge, 371

soft coral and crinoids seen on the debris avalanche deposit blocks is not present. There are 372

sparse, scattered lithic angular clasts a few tens of centimeters across, in the floors of some of the 373

grooves, very similar to the clasts forming the blocks of the matrix rich debris avalanche facies . 374

Apart from these clasts, the groove floors are largely covered by soft sediment. The best 375

exposures of grooves are on top of the sediment mesas, although they are also exposed in the 376

intervening areas. Isolated mound features, several tens of meters across and with high 377

backscatter intensity are also present in the grooved terrain. 378

379

5.3. Distal debris avalanche flow facies 380

381

The transition between the matrix rich debris avalanche facies and the distal deposits to 382

the north is gradual without a sharp boundary that can be differentiated in the seismic data. The 383

hummocks that characterized the matrix rich facies with presence of debris avalanche blocks 384

decrease to leave a relatively flat seabed of acoustically transparent material in the distal deposits. 385

This type of transition is common in debris avalanches that transform to debris flows, whilst in 386

debris avalanche lobes where this transformation does not occur the lobe has an abrupt front 387

(Capra et al., 2001). We mapped an approximate limit for this transformation, based on the 388

difference in backscatter intensity. The distinction between the two facies is also supported by 389

comparison of TowCam images from tows 5 and 6 (matrix facies) with tows 3 and 4 (distal 390

facies). Here we discuss the depositional and erosional features of the distal facies. 391

392

5.3.1 Depositional features 393

394

The distal landslide deposit consists of a laterally extensive sheet extending over a broad, 395

almost flat submarine plain with an average slope of 0.11º. Within the distal deposit are areas with 396

differential high or low backscatter intensity that correspond to variations on the slope but that do 397

not show acoustic variations in the CHIRP profiles. The margins of the deposit are ill-defined and 398

in at least some places covered by more recent sediment (see below). It is likely that the distal 399

deposit extends out of the northern limit of backscatter mapping beneath the “undivided marine 400

sediments” unit of Figure 12. Figure 11 shows representative portions of seismic lines along and 401

across the distal facies, where the deposit is recognized as a transparent unit beneath a strong echo 402

and overlying a sedimentary sequence. In contrast to the matrix-rich debris avalanche facies, the 403

distal facies lacks the characteristic hummock facies typical of debris avalanches. We used a 404

nominal water velocity of 1500 m/s to estimate the thickness of the deposits. The minimum 405

thickness of the deposit is about 2 m in areas with somewhat greater slope. The deposit is 5-10 m 406

thick in the areas where we recognized a reflection beneath the transparent acoustic facies, but in 407

many areas it may be thicker, because the CHIRP system failed to image any reflector beneath the 408

avalanche. 409

The TowCam images across the distal deposits show that the seabed exposes areas of soft 410

sediment, as indicated by the presence of bioturbation and the lack of sessile epifauna, and areas 411

with abundant dark, rounded clasts 10-20 cm across partially covered by lighter sediment (Figure 412

10h). In the transitions from areas with low to high clast abundance, we observed clast alignments 413

and concentrations in ridges. The clast alignments are oriented SW to NE on the seabed 414

(TowCam 3, Figure 3) parallel to a fabric seen in the backscatter intensity data. 415

Dredge samples indicate that the clasts are cohesive silt-grade sediment (Figure 10i, see 416

TowCam 3 in Figure 3 for location). The size of the clasts in the dredge is comparable to the size 417

of the dark clasts in the camera images. Comparison with the sizes of the scallops seen in the 418

grooved terrain and an inferred similarity between the silt-grade sediments in the clasts found in 419

the dredge and the cohesive sediments seen in the grooved terrain and the walls of the sediment 420

mesas suggests that the debris flow deposits contain a high proportion of sediments eroded by the 421

landslide. In contrast, only four small lithic clasts were present in dredge sample: a reworked 422

dacitic pumice pebble, two clasts of andesitic and dacitic composition and a basaltic fragment of 423

Mg-rich olivine and chrome spinel–phyric; not being the latter similar to the basalts of any of the 424

adjacent volcanic islands (Johnson et al., 1972). 425

In the northwestern margin of the distal deposit we mapped a series of arcuate features in 426

the backscatter that were recognized as sediment waves coming from the island of Tolokiwa 427

(Hoffman et al., 2008) (Figures 3 and 12) and these are also seen in the chirp data. We also 428

interpret the existence of sediment waves surrounding the eastern margin of the landslide deposit, 429

coming in this case from the Dampier strait. 430

431


433

No erosional features were observed over the main area of the distal deposit. However, its 434

southern margin extends along the foot of the line of sediment mesas discussed in the previous 435

section, whilst its western margin is located at the foot of a scarp some tens of metres high eroded 436

into marine sediments that was partly imaged in TowCam 4 (Figure 3) crossing that margin. 437

Scalloped features at the crest of the scarp, similar to those seen in the grooved terrain, were 438

imaged by the TowCam (Figure 10x and y). It is therefore possible that the southern part of the 439

distal deposit (X in Figure X. SIMON, I need the exact location for this area), covering an area 440

some 20 km (E-W) by 10 km (N-S) lies within an area where mass erosion or failure of sediment 441

in the substrate excavated the seafloor to a depth of up to 20 to 40 m (the height of the highest 442

sediment mesas above their surroundings). 443

444

6. Volume estimates 445

446

6.1. Volume of the scar 447

Volume estimation of a volcanic island flank collapse where a well-defined escarpment or 448

amphitheater exists can be done by subtracting a grid of the present topography from a grid of the 449

calculated pre-collapse surface. The latter is usually obtained by extrapolating the topographic 450

and bathymetric contours across the embayment (e.g., Urgeles et al, 1997). For the Ritter 451

collapse, where most of the subaerial part of the island and the inferred convex submarine flank 452

has disappeared in a single event, it is necessary not just to extrapolate, but also to restore the 453

whole original island’s topography. 454

Restoration of the pre-collapse topography of Ritter together with an extrapolation of the 455

submarine collapse scar as a single curved surface approximates the initial pre-collapsed shape of 456

the island and its submarine failed flank (Figure 13a). To create the pre-collapse grid it was 457

necessary to remove the landslide blocks, the post-collapse volcanic eruptions and to smooth the 458

erosion of the submarine valley assumed to have been excavated during the event. Subtracting the 459

reconstructed slide surface (Figure 13b) from the reconstructed pre-collapse topography we 460

obtained a total volume of 4.2 km3 for the slide mass involved in the collapse of Ritter. The 461

contoured difference between these two surfaces used to calculate the volume is shown in Figure 462

13c. 463

464

6.2. Volume of the deposits 465

466

In comparing collapse scar and deposit volumes, it is important to recognize the 467

occurrence of dilation during failure and landslide movement. The deposits of the well 468

constrained 1980 Mount St. Helens debris avalanche were estimated to have experienced around 469

25% dilation over their pre-collapse volume (Glicken; 1996). Gridding the thickness of the block 470

facies in the avalanche within the limits of the Ritter deposit mapped provided a volume of 4 km3. 471

Nevertheless, if we correct the latter value for a 25 % of dilation during the collapse, a volume of 472

3 km3 of the initial collapse material would have been necessary to form the large block debris 473

avalanche facies. Because we have obtained a volume of 4.2 km3 for the Ritter Island subaerial 474

and submarine scar, we infer that a fraction of about 71% of the initial landslide material never 475

got beyond the constriction in the valley between Sakar and Umboi islands, and is now emplaced 476

as the large block facies deposit. Only a smaller volume of material reached farther downslope 477

and was highly efficient in mobilizing the substrate. 478

Estimates of the volume of the matrix rich debris avalanche facies deposit and of the 479

volume of material eroded by the passage of the landslide before its emplacement are uncertain 480

because none of the chirp lines provided deposit thickness in this area. We infer that the bulk of 481

the distal debris flow deposit is largely not from the initial collapse but rather from entrained 482

sediment, which is consistent with the low abundance (<1%) of lithic clasts in the dredge sample. 483

Estimation of the volume of the distal deposit is complicated by the sediment cover on the eastern 484

margin and by its uncertain northern limit. It is likely that the distal deposit extends out of the 485

northern limit of backscatter mapping (Figure 3). Within the mapped region the deposit area is 486

about 800 km2. Silver et al. (2009) estimated an average thickness of 46 m for the block-rich part 487

of the debris avalanche deposit. For the Canary Island collapses, Masson et al. (2002) reported 488

average thicknesses of 100 m (El Golfo slide), 10 m (Canary slide), 23 m (Saharan slide), 75 m 489

(Guimar slide), and 122 m (Cumbre Nueva slide); and Llanes et al. (2003) reported an average 490

thickness of 72 m for Anaga debris avalanche. Satake and Kato (2001) reported an average 491

thickness of 36 m for the Oshima-Oshima debris avalanche; Francis et al. (1985) estimated 492

average thickness of the Socompa deposit as 31 m; Deplus et al. (2001) measured 225 m near the 493

distal edge of the St. Lucia debris avalanche, and Belousov and Belousova (1998) reported 494

average thicknesses of 30 m, 14 m, and 9 m, each for the northern, central and southern parts of 495

the March 30, 1956 Kezymyannyi (Kamchatka) debris avalanche deposits. 496

The thickness of the Ritter debris avalanche deposit varies within the mapped region. High 497

frequency CHIRP seismic data taken over the distal facies (Figure 11) show thicknesses of 7-10 498

m for much of the region for which we have data. If we assume a conservative average thickness 499

of 8 m, the minimum volume of the avalanche should be 6.4 km3. This conservative value would 500

imply that at most 15% of the distal deposit is composed of lithic clasts originating from the 501

initial collapse of the Ritter Island edifice, and likely less because of the finite although unknown 502

volume of the matrix-rich debris avalanche facies deposit. 503

504

7. Discussion 505

506

7.1. Instability and triggering of edifice collapse 507

508

Our data do not provide direct evidence on the processes that triggered the March 13th 509

1888 lateral collapse of Ritter Island, but they they provide potential controls on the geometry of 510

the collapse and hence factors that may have made the edifice susceptible to failure. 511

The 1888 Ritter Island sector collapse differs from most subaerial volcano collapses in 512

that the collapse scar cuts down to the base of the edifice and perhaps into the substrate. Subaerial 513

lateral collapses typically cut half way down to the base of the pre-collapse edifice, like the 1980 514

Mount St. Helens collapse (Siebert et al., 1987). In a few subaerial volcanoes and other submarine 515

or island volcanoes, the failure occurs at the base of the edifice or in the substrate. At the 516

Socompa volcano, it has been argued that the reason for failure at the base of the volcano lies in 517

pre-collapse deformation of weak, water-saturated sediments at the base of the edifice (Francis et 518

al., 1985; Van Wyk de Vries et al., 2001). This deformation is manifested in a series of anticlines 519

extending beyond the foot of the volcano, the crests of which form ridges that now project 520

through the landslide deposit (Van Wyk de Vries et al., 2001). In the data from Ritter Island, we 521

did not find any evidence of pre-collapse deformation. The Ritter volcano was most likely built 522

upon marine sediments that could have provided a weak basal layer that localized the failure at 523

the base of the volcano. Other examples of volcanoes cutting down to the base of the edifice are 524

the failure of Unzen in 1792 (Siebert et al. 1987) and the collapse scar at Bandai, that cuts 525

unusually deep into the edifice (Siebert et al. 1987). Both of these lateral collapses were 526

associated with phreatic hydrothermal explosions and hot groundwater discharges, and lacked 527

associated large-scale magmatic eruptions, similar to Ritter. 528

There is no evidence that tectonic triggers such as large-scale fault movements or 529

gravitational settling-related faulting initiated the Ritter landslide. Environmental factors such as 530

high-intensity rainstorms, which can cause a reduction in strength of the edifice due to saturation 531

(McGuire, 2003), were also absent. The collapse lacked an associated large-volume magmatic 532

eruption, but was likely associated with phreatic or phreatomagmatic explosions, because 533

observations of vapor emission at Ritter were reported only a week before the collapse (Cooke, 534

1981), indicating that a shallow hydrothermal system was present within the edifice (Siebert et al., 535

1987). The Bandai and Unzen types of lateral collapse may be triggered by pore fluid 536

pressurization deep within the edifice due to fluid input and / or heating from magmatic intrusions 537

emplaced well below the base of the edifice, rather than within the edifice as in the case of the 538

1980 Mount St. Helens cryptodome (Day, 1996; Reid, 2004). In these cases, the locus of 539

pressurization, deformation and failure may be in low permeability layers near or at the base of 540

the volcano: this could be another mechanism by which a substrate of compacted deep marine 541

sediments could account for the geometry of the 1888 Ritter Island collapse structure. 542

There is no geological evidence in our data for multiple or retrogressive failures of the 543

Ritter Island volcano within the 1888 collapse event, but we cannot exclude the possibility that 544

lateral collapse at Ritter involved successive blocks beginning to move tens of seconds apart. 545

Retrogressive failures with blocks beginning to move a few tens of seconds apart and successive 546

failures a few minutes apart have been directly observed by eyewitnesses at Mount St. Helens in 547

1980 (Voight, 1981) and Stromboli in 2002 (Bonaccorso et al., 2003) respectively, but the effects 548

in the deposits are subtle and ambiguous (Glicken, 1996; Ward & Day, 2006). Evidence that the 549

collapse at Ritter Island began as a single failure or as a rapid sequence of retrogressive failures 550

no more than a few tens of seconds apart, is principally provided by the single tsunami that it 551

produced (Cooke, 1981; Ward & Day, 2003). 552

553

7.2. Landslide erosion 554

555

7.2.1 Scouring 556

557

The scouring mark found in the slope of Umboi island (Figure 5) lacks a distinct 558

bathymetric expression and the upslope boundary of the very high backscatter zone has no 559

expression in the chirp profile that crosses it, unlike the boundaries of the other deposits that we 560

associate with the 1888 Ritter Island collapse in this paper. Similarly, it appears unlikely that it is 561

a deposit originating from further upslope, since all other such deposits evident in the backscatter 562

data from Umboi and also Sakar trend down the slope. Furthermore, it is not an artifact due to the 563

angle relative to the beam, because in that case the upslope of the contact would then have the 564

higher backscatter not the lower one. Instead, in the light of the undulating nature of the upslope 565

contact and the way it passes upslope over of the small ridge, an alternative interpretation is that it 566

is an erosional feature comparable to the scouring of the walls of Coldwater creek by a fast-567

moving lobe of the 1980 Mount St. Helens debris avalanche (Voight et al., 1981; Fisher et al., 568

1987; Glicken 1996). Figure 5 compares the two features. 569

The landslide moving at high velocity, obliquely impacting the slope of Umboi and 570

running up the slope, rising as it impinged on the small ridge, may explain the pattern of inferred 571

erosion. This feature is not simply a “high tide” mark recording the maximum thickness of the 572

landslide as it moved parallel to the slope, because this does not account for the sharp rise in the 573

upslope contact up and over the small ridge, nor does it explain why a comparable feature is not 574

observed at the same elevations on the slope of Sakar. The orientation of the collapse scar, which 575

faces obliquely towards Umboi and away from Sakar may explain the lack of a corresponding 576

feature on the slopes of Sakar. 577

578

7.2.2. Black Streak 579

580

The very high intensity backscatter character of this unit, together with its differential 581

relief in the bathymetry and the uniformity of the size clasts imaged with the photographs, lead us 582

to interpret the Black Streak as a volcanic unit. The uniformity of the size clasts (10 to 20 cm 583

across) is rarely seen in debris avalanche deposits, but is typical of scoriaceous volcanic units. 584

Thus, we infer that this feature is the core of an in-situ submarine volcanic vent whose top was 585

eroded off by the leading part of the landslide. 586

587

7.2.3. Mesas 588

589

The development of scarps and troughs in the mesas located to either side of the lobes of 590

the matrix-rich debris avalanche deposit (M in Figure 9) lead us to interpret them as erosional 591

features related to a flow moving from SE to NW, similar to the groove trend seen to the SE. The 592

formation of the mesas must predate emplacement of the matrix-rich debris avalanche facies lobes 593

in the troughs between the mesas, but the association between them indicates that they are both 594

related to the Ritter landslide event. However, the mesas and troughs located along the SW 595

margin of the matrix rich facies deposit (M’ in Fig. 9, Line 3/1 in Fig. 10) may not be originated 596

by a flow moving down-slope, but instead they may be the result of uneven upslope retreat of 597

scarps developed at the lateral margin of the moving debris avalanche, as a result of undermining 598

the foot of the slope by erosion along the valley axis, followed by development of progressive 599

slumping upslope. 600

601

7.2.4. Grooved terrain 602

603

The groove walls have strength sufficient for the delicate sponges to attach themselves, 604

but are too weak to support the other elements of the fauna seen on the debris avalanche blocks. 605

Scalloped but linear grooves eroded in cohesive water-saturated sediment have been produced 606

experimentally by high velocity water flows simulating tidal flows over littoral muds 607

(Winterwerp & Van Kesteren, 2004). The grooves imaged in our work are very similar to the 608

intermediate-stress bulk erosion regime, as characterize experimentally by the latter work. We 609

infer that the grooved terrain was generated by bulk erosion of older marine sediments at the base 610

of a high velocity component of the Ritter debris avalanche that extended over a wider front than 611

the narrow matrix-rich debris avalanche deposit. Further evidence for the association of the 612

grooves with the Ritter landslide is provided by the occurrence of sparse, scattered lithic angular 613

clasts a few tens of centimeters across, in the floors of some of the grooves, very similar to the 614

clasts forming the blocks of the matrix rich debris avalanche facies (Figure 10). Isolated mound 615

features, several tens of meters across and with high backscatter intensity, are also present in the 616

grooved terrain, which we interpret as isolated debris avalanche blocks. 617

618

7.3. Inferred landslide kinematics 619

620

The collapse of Ritter left a deposit composed of three different facies, a division that 621

implies three distinct stages in the movement and rheological evolution of the landslide. The 622

distribution of these facies and the different types of erosional structure, together with the overall 623

geometry of the deposit can be used to reconstruct the movement and evolution of the landslide, 624

as Glicken (1991; 1996) showed could be done in the case of the 1980 Mount St. Helens collapse. 625

The grooved terrain and the sediment mesas indicate the occurrence of extensive erosion, 626

of at least the top 20 meters of the existing sediment sequence in these areas. The association of 627

the grooved terrain with isolated debris avalanche blocks, the occurrence of scattered lithic clasts 628

in the smaller grooves imaged with the TowCam, and the channelization of the debris avalanche 629

lobes between the sediment mesas, indicate that the erosion took place during the landslide, and 630

that deposition of the matrix rich facies debris avalanche deposit took place after the substrate 631

erosion. We conclude that the erosional features were produced by the fast moving leading part of 632

the landslide that, after being deflected off the submarine slope of Umboi in a “bobsled” fashion 633

similar to that seen in some subaerial landslides in mountain valleys (Evans et al., 1989), reached 634

beyond the constriction and spread out over a front up to 10 km wide (corresponding to the width 635

of the grooved terrain and the line of sediment mesas). The distribution of the grooved terrain and 636

the sediment mesas indicate that it was this part of the landslide that generated the substrate 637

erosion and the debris flow. In contrast the matrix rich debris avalanche was deposited from the 638

slower, trailing part that was confined to the axis of the valley between Umboi and Sakar. 639

The landslide pushed waves up that devastated the coast of Sakar and Umboi islands to 640

more than 15 m of elevation (Ward and Day, 2003). Run-up of landslides onto valley walls can be 641

also used to estimate landslide velocities, by equating the potential energy gained in the run-up to 642

the minimum kinetic energy of the moving material at the base of the slope. This argument was 643

applied to landslides by Heim (1882) and had notable applications to the debris avalanche of 644

Mount St. Helens (Voight et al., 1983) and to the Osceola mudflow at Rainier (Vallance & Scott, 645

1997). In Ritter, some landslide material ran up and over the slopes of Umboi and Sakar islands, 646

implying that it had sufficient kinetic energy to climb uphill. Assuming that the general level of 647

the landslide was at the 900 m level and applying this criterion would suggest that the material 648

that climbed up to 200 m over the small ridge (Figure 4) would have had a speed of order 63 m/s. 649

This is actually a high estimate because of the fact that it neglects the particle-particle interaction 650

within the debris avalanche that would tend to push particles up the slope. Ward and Day (2003) 651

obtained a more conservative velocity of 40 m/s for Ritter landslide taking into account particle 652

interaction. However, this rate is not necessarily that of the landslide as a whole, since it is likely 653

to have fragmented much earlier in its motion and the front may have been traveling significantly 654

faster than the tail. 655

656

7.4. Transformation from debris avalanche to debris flow 657

658

The SW-NE orientation of clast alignments seen on the seabed of the distal facies 659

(TowCam line 3, Figure 3) parallel to a fabric seen in the backscatter intensity data, probably 660

follows the landslide movement direction in this area and indicates pervasive shearing during 661

deposition (Figure 10g). Such structure is not seen in debris avalanches, where shear is 662

concentrated at the base and margins producing marginal levees, nor in turbidite deposits, but are 663

often seen in debris flows (REFERENCES NEEDED HERE). On account of this clast alignment 664

fabric, the low slope of emplacement and inferred very low aspect ratio of the deposit, and the 665

lack of hummocks, we interpret the distal deposit as a debris flow unit. 666

Previous work has shown that progressive disintegration of landslide blocks can generate 667

debris flow deposits extending downslope from the original landslide (Piper et al., 1999; Gee et 668

al., 1999; Scott et al., 2001). These transformations often occur by erosion and entrainment of 669

water rich sediments, glacier ice and other debris from the substrates over which the landslides 670

move. Where there is abundant substrate that is easily entrained, as in the floors of glaciated 671

valleys, the total volume of both volcanic and non volcanic landslides can increase by more than 672

1000% over the initial failure volume (Plafker & Ericksen, 1978; Scott et al., 2001; Erismann & 673

Abele, 2001; Kaab et al., 2003), a process known as bulking (Scott et al., 2001). Bulking 674

increases the total landslide runout distance and the landslide mobility, given by the decrease in 675

drop height / run-out distance ratio and commonly interpreted as effective friction coefficient. 676

Understanding the development of landslide facies and the role of bulking is therefore key to the 677

understanding of landslide mechanics and potential tsunamis. 678

The transformation of the leading part of the Ritter debris avalanche into the debris flow 679

was also related to substrate erosion and entrainment, as suggested by the dominance of sediment 680

intraclasts in the dredge sample and in the TowCam images, in contrast to the lithic clast 681

dominated mounds seen in the matrix rich debris avalanche facies. The exposed areas of substrate 682

erosion in the grooved terrain and the sediment mesas upslope of the present transition from 683

debris avalanche deposit to debris flow deposit are likely sources of the entrained sediment. In 684

addition, the scarps along the western margin of the deposit indicate that another source is the 20 685

km by 10 km area bounded by the scarp to the west and the sediment mesas to the southeast and 686

now covered by the upslope, proximal parts of the debris flow deposit itself (SIMON, I need the 687

exact location of this area in Figure 12). The lack of distinct raised sediment mesas within the 688

latter area implies that such erosion must have involved complete removal of the substrate by 689

mass failure, the most complete mechanism of substrate erosion identified by Winterwerp & Van 690

Kesteren (2004). The coincidence of this change from mass erosion in the grooved terrain, to 691

mass failure around and to the N of the sediment mesas, with the margin of the near-flat plain on 692

which the debris flow was deposited, may be due to a transition from stronger or coarser sediment 693

on the submarine slopes of Umboi and Sakar to weaker, finer grained sediment on the plain that 694

was more susceptible to failure under the impact of the fast-moving leading part of the landslide. 695

Alternatively, Talling et al. (2007) showed how debris flows could deposit very little 696

sediment and be locally very erosive in large areas far from the sources, and then sediment 697

deposition could be triggered by a remarkably small but abrupt decrease in sea-floor gradient 698

from 0.05° to 0.01°. In this case, our data show that there is an order of magnitude reduction in 699

slope between the area of debris avalanche (1.6-2.1º) and the area of debris flow deposition (0.1º). 700

The coincidence of the facies transition and the break in slope could indicate that the debris-flow 701

deposition started due to rapid deceleration of the flow, because deceleration reduces the turbulent 702

mixing and sand settles more rapidly, as interpreted by Talling et al. (2007); however, this process 703

does not explain the onset of mass failure at the break in slope. The two mechanisms may not be 704

mutually exclusive, because conservation of momentum requires that a sudden bulking of the 705

landslide due to mass failure of the substrate would be accompanied by a reduction in its velocity. 706

Debris flow and turbidite deposits associated with 1888 Ritter collapse may be found 707

much farther to the north and west in the Bismarck Sea Basin than the area mapped, similarly to 708

what has been identified associated with the 1929 Grand Banks continental slope failure (Piper et 709

al., 1985) and with the Canary (Talling et al., 2007; Hunt et al., 2011) and Hawaii (Kanamatsu et 710

al., 2002) landslides. 711

712

CONCLUSIONS 713

714

1) The collapse of Ritter Island in 1888 removed most of the island and was probably 715

associated with phreatic or phreatomagmatic explosions. Based on our mapping, the collapse scar 716

cut down into the base of the edifice. Most subaerial volcano collapses cut only part-way to the 717

edifice base. We conclude that the collapse resulted from a single event because in mapping the 718

full seafloor around the island we found no field evidence either onshore or offshore that more 719

than one collapse occurred then or in the past. Historical accounts and tsunami modeling also are 720

consistent with a single event. 721

2) The deposits of the landslide generated by the collapse comprise three differentiated 722

facies: large block, matrix-rich and debris flow distal facies. The transitions from large block to 723

matrix-rich facies and from the latter to distal facies are related to major changes in the slopes and 724

constrictions. The facies are well differentiated by their characteristics, as observed in our various 725

datasets including bathymetry, side-scan sonar, high-resolution seismic data and tow-cam 726

photographs. 727

3) Tow-cam photographs of the sea-bottom provided important constraints on the remote 728

sensing and are very useful for studying landslide facies in subdued areas. Characteristics seen in 729

the photos provided valuable information for mapping the facies of the avalanche including: 730

morphology of the deposit, clasts aspect, distribution and abundance, presence/absence and type 731

of organisms indicative of substrate strength, and signs of erosion such as scalloped walls. 732

4) Volumetric calculations based on the restoration of a modeled original topography led 733

to a total volume estimate of 4.2 km3 for the initial Ritter Island collapse. After correcting from 734

dilatation, we infer that about 75% of the material from the initial collapse now forms the large 735

block facies at distances of less than 12 km from the collapse scar, filling an area between the 736

islands of Umboi and Sakar and topographically constricted by volcanic ridges from those 737

islands. The volume of the matrix-rich facies is unknown, but the distal debris flow facies 738

deposits have a minimum volume of 6.4 km3, as a result of bulking of the landslide through 739

sediment entrainment. This highlights the efficiency of substrate erosion during the later history 740

of the landslide movement, when much of its original volume had been retained in the large block 741

facies deposits upslope. 742

5) Recent submarine landslide deposits that are well-exposed can be interpreted similarly 743

to subaerial volcano collapse deposits. In turn, these deposits can provide a basis for the 744

interpretation of older and incompletely exposed landslide deposits. Understanding the 745

mechanisms and kinematics of collapse events can improve tsunami modeling. 746

747

ACKNOWLEDGMENTS 748

749

This work was supported by National Science Foundation grant OCE-0327004 to Eli 750

Silver and Steve Ward, and grant OCE-0328278 to Neal Driscoll. Participation of Pilar Llanes 751

was possible thanks to grants “Becas del Amo 2008” from Universidad Complutense de Madrid 752

and “Movilidad para Profesores 2011” from Caja Madrid Foundation. We are grateful for the 753

advice and support of many people in carrying out this research, including Dan Fornari, Bruce 754

Appelgate, Jennifer Engels, Steve Saunders, Hugh Davies, James Robins, and Amilia Lyons. We 755

are indebted to the officers, crew, and scientific party of the R/V Kilo Moana for their cooperation 756

at sea. Figs. 1, 3, 4, 5, and 13 were drawn using GMT software (Wessel and Smith, 1995). 757

758

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study from Icod landslide on north Tenerife, Canary Islands. Geochem. Geophys. Geosyst. 835

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Avalanches. 1: Natural Phenomena, ed. B Voight, pp. 277–314. Amsterdam: Elsevier 863

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debris flows transformed from landslides in volcanic terrains: mobility, hazard assessment 876

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avalanche of May 18. In: Lipman, P.W. & Mullineaux, D.R. (eds.), 1981. The 1980 900

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Mount St. Helens rockslide-avalanche of 18 May 1980. Geotechnique 33, 243-273. 904

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interpretation of deposits and landslide seismic signals of Mount Saint Helens, 1980 May 908

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Islands. Journal of Geophysical Research 100, 24487 – 24498. 911

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http:/www.agu.org/eos_els, American Geophysical Union. 913

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Sediment in the Marine Environment. Developments in Sedimentology 56, 466 pp. 915

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geochemistry of the 30° - 35° S sector, and overview, of southern Kermadec arc 917

volcanism. Journal of Volcanology and Geothermal Research 149, 263 – 296. 918

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Hawaii Island: sampling and observations by remotely-operated vehicle Kaiko. Bull. Of 920

Vol. 66, 569-584. 921

922

FIGURE CAPTIONS 923

924

Figure 1. Location of Umboi, Sakar and Ritter islands over digital terrain model (DTM) 925

constructed from multibeam data gridded at 24 m. Topography from SRTM data. Inset shows 926

location in the New Britain volcanic arc, northeast of Papua New Guinea. 6’ of latitude = 11.1 927

km. 928

929

Figure 2. (Top) Photographs of Ritter taken in 2006, courtesy of John Holder. Left: view from the 930

west. Right: view from the southeast. (Bottom left) Sketch of Ritter Island in 1835, modified from 931

Jacobs (1844). Arrows show how the modern island was used to estimate the scale of the sketch. 932

(Bottom right) Sketch of Ritter Island in 1835, modified from Woodworth (unpublished, courtesy 933

of James Fairhead, University of Sussex, UK). 934

935

Figure 3. Side-scan imagery of study area from MR-1 sonar data gridded at 16 m. White dotted 936

line outlines landslide deposits. BS marks location of the black streak. Yellow lines indicate chirp 937

seismic reflection profiles acquired and red lines indicate segments of those shown in the paper. 938

Transects using the underwater digital towed camera system “TowCam” are marked with dashed 939

blue line. Blue stars show location of core sites (PC for piston core and GC for gravity core) and 940

the blue circle of a dredge site. 941

942

Figure 4. Bathymetry map of study area gridded at 24 m. Contour interval is 100 m. Bathymetric 943

ridges from Umboi and Sakar constricting the valley are marked with R1 and R2 respectively. R3 944

marks the smaller ridge that the fast-moving part of the avalanche passed over. 6’ of latitude = 945

11.1 km. 946

947

Figure 5. Side-scan image detail of the large block debris avalanche facies from MR-1 sonar data 948

gridded at 16 m. Largest block are marked as B and post-collapse volcanics as V. Areas of 949

sediment scouring are labeled (see also Fig. 12, “scouring to bedrock by landslide” in inset 950

legend). R1, R2 and R3 mark bathymetric ridges constricting the valley, the later being the 951

smaller ridge that the fast-moving part of the avalanche passed over. Low backscatter light and 952

high backscatter dark. Topography from SRTM data. 6’ of latitude = 11.1 km. 953

954

Figure 6. 3D image of Ritter insular margin. Letter B marks location of the largest debris 955

avalanche block. Note how the large block facies has been confined near the island by a 956

constriction. Letter V marks location of the main post-collapse volcanic cone. Vertical 957

exaggeration is x 2. 958

959

Figure 7. Part of CHIRP seismic line 7/0 across the large block avalanche facies. The collapse 960

blocks produced strong seabed diffractions and lack internal structures. Location of the lines is 961

shown in Figure 3. TWT denotes two way time. 962

963

Figure 8. Side-scan image detail of matrix-rich debris avalanche facies from MR-1 sonar data 964

gridded at 16 m. Low backscatter light and high backscatter dark. M for sediment mesas with 965

marked troughs around their margins in some cases. We interpret the mesas to be erosional 966

features related to a landslide or flow moving from SE to W. M* for the best example of a 967

sediment mesa (see text for details). Figure 11 shows a chirp profile crossing two of the mesas, 968

included M*. M´ for sediment mesas along the SW margin of the facies deposit and without 969

troughs developed around them. We attribute the latter to be the result of uneven upslope retreat 970

of scarps. 971

972

Figure 9. (Top) Chirp line 2/7 showing matrix-rich facies of Ritter debris avalanche and its 973

transition towards the east to a layered sedimentation at the lower insular slope of Sakar island. 974

(Middle) Continuation of Chirp Line 2/7 showing large blocks and a buried landslide deposit 975

from Sakar island. (Bottom) Line 3/1 across the insular slopes of Umboi and Sakar island 976

showing the matrix-rich facies of Ritter and the channel eroded by Ritter debris avalanche. 977

Diffractions related to hummocks. Location of the lines is shown in Figure 3. TWT denotes two 978

way time. 979

980

Figure 10. Photograph images taken with the “TowCam” system. A) Rock clasts from a 981

hummock within the matrix-rich debris avalanche facies. B) Polygonal-jointed columnar blocks 982

in one of the hummocks within the matrix-rich debris avalanche facies. The polygonal jointing 983

suggest either dike rocks or part of a thick subaerial lava flow. C) Steep side of a hummock 984

composed of pitted clasts appartently coarsely vesicular. D) Angular clasts, dark and pitted from 985

the Black Streak. The clasts are uniform in size, around 10-20 cm across, in contrast to the very 986

poorly sorted clasts of the debris avalanche hummocks, suggesting sorting in a submarine 987

explosive eruption. E) Grooved terrain with scalloped walls. The wall is clearly overhanging, 988

indicating that the sediments are strongly cohesive F) Grooved terrain cutting into cohesive 989

sediments and with lithic angular clasts in the floor down the wall. G) Clast alignment seen in the 990

distal facies following the landslide movement direction and indicative of pervasive shearing 991

during deposition. H) Cohesive sediment clasts seen in the distal facies. I) Same clasts as seen in 992

H taken from the dredge sample. 993

994

Figure 11. Segment of CHIRP LINE 24 across the matrix-rich debris avalanche showing the pre-995

avalanche sediments, the “black streak” and sediment mesas. Lines 3/25, 2/4 and 2/5 are selected 996

segments of CHIRP profiles across the distal facies, while Line 22 shows a segment across it. Red 997

bars mark thickness of the avalanche, labeled in meters using a conservative velocity of 1500 m/s 998

to calculate the thickness (if a velocity of 1600 m/s will be used, the values will increased by a 999

6.25%). Location of the lines is shown in Figure 3. TWT denotes two way time. 1000

1001

Figure 12. Morphological interpretation of Ritter insular margin and adjacent areas. See inserted 1002

legend for interpretation of the different morphological features. 1003

1004

Figure 13. A) Topographic and bathymetric reconstruction of Ritter Island and its insular margin 1005

before its collapse on March 13th, 1888. B) Topography of Ritter Island and insular margin after 1006

its collapse but without the post-collapse volcanic vents. C) Thickness of displaced slide mass, 1007

obtained by subtracting the latter to the former reconstructions. 1008

, pilar llanes2,3*, eli silver2, gary hoffmann , steve ...ward/papers/ritter-deposits.pdf · 20...

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